Environmental Soil Chemistry - Chapter 2 pps

A coordinated polyhedron of anions is formed about each cation, the cation–anion distance being determined by the radius sum sum of cation and anionradii and the coordination number of t

Components

Introduction

Soils are complex assemblies of solids, liquids, and gases For example,

in a typical silt loam soil ideal for plant growth the solid component

in the surface horizon represents about 50% of the volume (about45% mineral and 5% organic matter), gases (air) comprise about 20–30%,and water typically makes up the remaining 20–30% Of course, thedistribution of gases and water in the pore space component can changequickly depending on weather conditions and a host of other factors.The median and range of elemental content of soils from around theworld are given in Table 2.1 The elements that are found in the highestquantities are O, Si, Al, Fe, C, Ca, K, Na, and Mg These are also the majorelements found in the Earth’s crust and in sediments (Table 2.1) Oxygen isthe most prevalent element in the Earth’s crust and in soils It comprisesabout 47% of the Earth’s crust by weight and greater than 90% by volume(Berry and Mason, 1959)

The inorganic components of soils represent more than 90% of the solidcomponents Their properties such as size, surface area, and charge behavior

greatly affect many important equilibrium and kinetic reactions andprocesses in soils.

The inorganic components of soils include both primary and secondaryminerals (defined below), which range in size (particle diameter) from clay-sized colloids (<2 μm or 0.002 mm) to gravel (>2 mm) and rocks Table 2.2lists the major primary and secondary minerals that are found in soils

A mineral can be defined as a natural inorganic compound with definitephysical, chemical, and crystalline properties A primary mineral is one thathas not been altered chemically since its deposition and crystallization frommolten lava Examples of common primary minerals in soils include quartzand feldspar Other primary minerals found in soils in smaller quantitiesinclude pyroxenes, micas, amphiboles, and olivines Primary minerals occurprimarily in the sand (2–0.05 mm particle diameter) and silt (0.05–0.002 mmparticle diameter) fractions of soils but may be found in slightly weatheredclay-sized fractions A secondary mineral is one resulting from the weathering

of a primary mineral, either by an alteration in the structure or fromreprecipitation of the products of weathering (dissolution) of a primarymineral Common secondary minerals in soils are aluminosilicate mineralssuch as kaolinite and montmorillonite, oxides such as gibbsite, goethite, andbirnessite, amorphous materials such as imogolite and allophane, and sulfurand carbonate minerals The secondary minerals are primarily found in theclay fraction of the soil but can also be located in the silt fraction

Pauling’s Rules

Most of the mineral structures in soils are ionically bonded and theirstructures can be predicted based on Pauling’s Rules (Pauling, 1929) Ionicbonds are formed when an ion interacts with another ion of opposite charge

in the mineral structure to form a chemical bond Covalent bonds are thosewhich result from a sharing of electrons Most chemical bonds have acombination of ionic and covalent character For example, the Si–O bond isequally ionic and covalent The Al–O bond is approximately 40% covalentand 60% ionic (Sposito, 1989) Pauling’s Rules (1929) are provided belowwith a brief description of their meaning in soil mineral structures

A coordinated polyhedron of anions is formed about each cation, the cation–anion distance being determined by the radius sum (sum of cation and anionradii) and the coordination number of the cation by the radius ratio

Ionic radii (IR) of cations and anions commonly found in inorganic soilminerals and the coordination number and radius ratio (assuming oxygen is thedominant anion) of the common cations are given in Table 2.3 The coordinationnumber (CN), which is a function of the radius ratio, is the number of nearestanions surrounding the cation in a mineral In soils, cations in mineral structureshave coordination numbers of 4, 6, 8, or 12 The radius ratio is the ratio of the

RULE 1

TABLE 2.1. Contents of Elements in Soils, the Earth’s Crust, and Sediments

TABLE 2.1. Contents of Elements in Soils, the Earth’s Crust, and Sediments (contd)

bFrom Bowen (1979) and references therein, with permission Represents soil analyses from throughout the world.

cFrom Bowen (1979) and references therein, with permission.

TABLE 2.2. Common Primary and Secondary Minerals in Soils a

Primary Minerals

Muscovite KAl2(AlSi3O10) (OH)2

Biotite K(Mg,Fe)3(AlSi3O10) (OH)2

Olivine (Mg,Fe)2SiO4

Epidote Ca2(Al,Fe)3Si3O12(OH)

Tourmaline (Na,Ca) (Al,Fe 3+ , Li, Mg)3Al6(BO3)3(Si6O18) (OH)4

Montmorillonite M x(Al, Fe 2+ , Mg)4Si8O20(OH)4(M = interlayer metal cation)

Vermiculite (Al, Mg, Fe 3+ )4(Si, Al)8O20(OH)4

Chlorite [M Al (OH)6](Al, Mg)4(Si, Al)8O20(OH, F)4

Allophane Si3Al4O12·nH2O

Imogolite Si2Al4O10·5H2O

Goethite α-FeOOH

Hematite α-Fe O

TABLE 2.2. Common Primary and Secondary Minerals in Soils a (contd)

Jarosite KFe3(SO4)2(OH)6

aAdapted from “Mineralogy: Concepts, Descriptions, Determinations” by L G Berry and B Mason Copyright © 1959 by W H Freeman and Company Also adapted from C S Hurlbut, Jr., and C Klein, “Manual of Mineralogy,”19th ed Copyright © 1977 John Wiley & Sons, Inc Reprinted by permission of John Wiley & Sons, Inc.

bAn explanation for the chemical formula can be found in the text.

cFormulas are for the full-cell chemical formula unit.

cation radius to the anion radius Cations having ionic radii less than a criticalminimum radius ratio can fit between closely packed anions having differentconfigurations (Fig 2.1) For elements in the same group of the periodic table the

IR increase as the atomic number increases For positive ions of the same electronicstructure the IR decrease with increasing valence For elements such as Mn thatexist in multiple valence states, the IR decrease with increasing positive valence.Figure 2.1 shows the relationship of radius ratio, coordination number,and the geometrical arrangements of nearest anions around a central cation.The IR of the oxygen ion in minerals is assumed to be 0.140 nm Based onFig 2.1 one can see that the Si4+cation would occur in fourfold or tetrahedralcoordination with O2–(the radius ratio would be 0.039/0.140 = 0.279); i.e.,four oxygen anions can surround the cation and result in a tetrahedralcoordination configuration such as that shown in Fig 2.1 Aluminum (Al3+)could also occur in fourfold coordination with O2–since the radius ratio is0.051/0.140 = 0.364 In fact, Al3+ can occur in either four- or sixfoldcoordination, depending on the temperature during crystallization of themineral High temperatures cause low coordination numbers, i.e., fourfoldcoordination, while at low temperatures sixfold coordination is favored.Based on the information given in Fig 2.1, Fe2+(0.074 nm), Fe3+(0.064 nm),and Mg2+ (0.066 nm) would occur in octahedral coordination As notedabove, Al3+ could also occur in octahedral coordination The coordinationnumber of these cations is 6, so that six O groups arrange themselves aroundthe cation, as shown in Fig 2.1

In a stable coordination structure the total strength of the valency bondswhich reach an anion from all neighboring cations is equal to the charge ofthe anion

RULE 2

RULE 3

This rule is known as the Electrostatic Valency Principle This can

be expressed as s = Z/CN, where s is the electrostatic bond strength to each coordinated anion, Z is the valence of the cation, and CN is the

coordination number (Pauling, 1929) Thus, for Si in tetrahedral

coordination, the electrostatic bond strength would be Z(4+) divided byCN(4), which equals 1 For Al in octahedral coordination the electrostatic

bond strength would be Z(3+) divided by CN(6) or 0.5 If Al substitutes for

Si in the tetrahedral layer, the electrostatic bond strength would be Z(3+)divided by CN(4) or 0.75, not 1 On the other hand, if Mg2+substitutes for

Al3+in the octahedral layer then the electrostatic bond strength is 2+/6 or0.33, not 0.5

The existence of edges, and particularly of faces, common to the anionpolyhedra in a coordinated structure decreases its stability; this effect is largefor cations with high valency and small coordination number, and isespecially large when the radius ratio approaches the lower limit of stability

of the polyhedron

Rule 3 is a statement of Coulomb’s Law for cations and indicates thatthere are three ways for tetrahedra and octahedra polyhedra (Figs 2.2A and2.2B, respectively) to bond: point-to-point, the most stable configuration,edge-to-edge, and face-to-face, the least stable configuration (Fig 2.3) With

TABLE 2.3. Ionic Radius (IR), Radius Ratio, and Coordination Number (CN) of Common Cations and Anions in Inorganic Soil Minerals a

Radius Ratio CoordinationNumber

cubo-FIGURE 2.1. Relationship of radius ratio, coordination number, and geometrical arrangement of nearest anions around a central cation Adapted from Dennen (1960), with permission.

cations with high valence such as Si4+ the tetrahedra are bonded point and with cations of lower valence such as Al3+the octahedra are bondededge-to-edge Polyhedra are not bonded face-to-face

point-to-In a crystal containing different cations those of high valency and small coordination number tend not to share polyhedron elements with each other

This rule is saying that highly charged cations stay as far from each other

as possible to lessen their contribution to the crystal’s Coulomb energy(Pauling, 1929)

The number of essentially different kinds of constituents in a crystal tends to

be small

This is because all substances tend to form the lowest possible potentialenergy Many kinds of constituents would result in a complex structurecharacterized by strains which would cause a high potential energy andinstability The different kinds of constituents refer to crystallographicconfigurations, tetrahedra and octahedra

RULE 4

RULE 5

FIGURE 2.2. (A) Diagrammatic sketch showing (a) single SiO 4 tetrahedron and (b) sheet structure of tetrahedra arranged in a hexagonal network (B) Diagrammatic sketch showing (a) single octahedra unit and (b) sheet of octahedral units From R E Grim, “Clay Mineralogy.” Copyright © 1968 McGraw–Hill Reproduced with permission of

McGraw–Hill, Inc.

A

B

FIGURE 2.3. The sharing of a corner, an edge, and a face by a pair of tetrahedra and

by a pair of octahedra From “The Nature of the Chemical Bond” by L Pauling, 3rd ed Copyright © 1960 Cornell University Used by permission of the publisher, Cornell

University Press.

Primary Soil Minerals

Some of the most important and prevalent primary minerals in soils are the feldspars (Table 2.2) They are common in the sand and silt fractions

of soils and can also be found in the clay fraction They comprise 59.5, 30.0, and 11.5% by weight of igneous rock, shale, and sandstone,respectively Metamorphic rocks also contain feldspars (Huang, 1989) The

K feldspars are important sources of K in soils and often compose a major component of the mineral form of soil K (Sparks, 1987) Feldspars are anhydrous three-dimensional aluminosilicates of linked SiO4 andAlO4tetrahedra that contain cavities that can hold Ca2+, Na+, K+, or Ba2+

to maintain electroneutrality (Huang, 1989) Feldspars can be divided into two main groups, alkali feldspars, ranging in composition fromKAlSi3O8 to NaAlSi3O8, and the plagioclases, ranging from NaAlSi3O8

to CaAl2Si2O8.Olivines, pyroxenes, and amphiboles are known as accessory minerals insoils and are found in the heavy specific gravity fractions Pyroxenes andamphiboles are ferromagnesian minerals with single- and double-chainstructures, respectively, of linked silica tetrahedra They make up 16.8% byweight of igneous rocks (Huang, 1989) Olivines are green neosilicates inwhich Mg2+ and Fe2+ are octahedrally coordinated by O atoms They areprevalent in igneous rocks, are sources of soil micronutrients, and aregenerally present in quantities smaller than those of pyroxenes andamphiboles (Huang, 1989)

Secondary Soil Minerals

Phyllosilicates

Clay is a general term for inorganic material that is <2 mm in size, whereasclay mineral refers to a specific mineral that mainly occurs in the clay-sizedfraction of the soil (Moore and Reynolds, 1989) Without question, thesecondary clay minerals (phyllosilicates) in soils play a profound role inaffecting numerous soil chemical reactions and processes as we shall see inthis chapter and in following chapters

Clay minerals are assemblages of tetrahedral and octahedral sheets (Fig 2.4) The silica tetrahedral sheet is characterized by a number ofproperties The Si–O bond distance is about 0.162 nm, the O–O distance isabout 0.264 nm, the tetrahedra are arranged so that all tips are pointing inthe same direction and the bases of all tetrahedra are in the same plane, andtetrahedra are bonded point-to-point

The aluminum octahedral sheet has an O–O distance of 0.267 nm and theOH–OH distance is 0.294 nm Bonding of Al octahedra occurs via edges.When one tetrahedral sheet is bonded to one octahedral sheet a 1:1 claymineral results Thus, the full-cell chemical formula for an ideal 1:1 claywould be Si4IVAl4VIO10(OH)8, where the superscripts represent four- andsixfold coordination in the tetrahedral and octahedral sheets, respectively.When two tetrahedral sheets are coordinated to one octahedral sheet, a 2:1clay mineral results The ideal full-cell chemical formula for a 2:1 claymineral would be Si8IVAl4VIO20(OH)4, e.g., pyrophyllite Between the sheets(i.e., interlayer space) cations may be octahedrally coordinated withhydroxyls, such as chlorites, and they may be present as individual cations,which may or may not be hydrated, as in micas, vermiculites, and smectites.Isomorphous substitution, which occurs when the mineral forms, is the

“substitution of one atom by another of similar size in the crystal lattice

without disrupting the crystal structure of the mineral” (Glossary of SoilScience Terms, 1987) Thus the size of the cationic radius determines whichcations can substitute in the tetrahedral and octahedral sheets In thetetrahedral sheet Al3+usually substitutes for Si4+but P can also substitute Inthe octahedral sheet Fe2+, Fe3+, Mg2+, Ni2+, Zn2+, or Cu2+can substitute for

Al3+ Thus, a cation with coordination number of 4 could substitute for Si4+

in the tetrahedral sheet and a cation of coordination number 6 couldsubstitute for Al3+in the octahedral sheet As a result of this isomorphoussubstitution, a net negative charge associated with the 6 oxygens or hydroxyls

of the octahedrons and with the 4 oxygens of the tetrahedrons develops

As an example of this suppose that one Al3+substitutes for one Si4+inthe tetrahedral sheet of an ideal 1:1 clay, Si4IVAl4VIO10(OH)8 Aftersubstitution the clay now has the formula (Si3Al1)IVAl4VIO10(OH)8 The totalnegative charge is –28 and the total positive charge is +27 The net charge onthe clay is –1, which is balanced by the presence of cations near the outersurface of the 1:1 clay

Clays can be classified as dioctahedral or trioctahedral, depending on thenumber of cation positions in the octahedral sheet that are occupied (Table 2.4)

If two of the three positions are filled, then the clay is dioctahedral If allthree positions are filled, the clay is trioctahedral For example, if Al3+ ispresent in the octahedral sheet, only two-thirds of the cation positions arefilled (dioctahedral); for every 6 OH– anions, only two Al3+ satisfy theanionic charge and a formula of Al2(OH)6results When Mg2+is present, allthree cation sites are filled since Mg is divalent and three atoms of Mg2+would be necessary to satisfy the 6 OH–ions (trioctahedral) The formula forthis sheet would be Mg3(OH)6

Trioctahedral minerals, which often contain Mg2+, are found in areaswith drier climates, while dioctahedral clays, which usually contain Al3+inthe octahedral sheet, are found in wet climates Thus, in the United States,east of the Mississippi River the soils have predominately dioctahedralminerals, while the clay fraction of soils west of the Mississippi is dominated

by trioctahedral minerals

FIGURE 2.4. Phyllosilicate nomenclature From Schulze (1989), with permission.

1:1 Clays (Kaolin–Serpentine Group). This group can be divided intodioctahedral kaolins and trioctahedral serpentines The most commondioctahedral kaolin is kaolinite (Table 2.4, Fig 2.5) Looking at the structure

of kaolinite (Fig 2.5), one sees that the basic structure consists of a silicatetrahedral sheet bonded to an aluminum octahedral sheet Layers ofkaolinite stack by hydrogen bonding (an electrostatic bond between apositively charged H+ion and a negatively charged ion such as O2–; Kleinand Hurlbut, 1985) and thus there are no interlayer spaces present, unlikewhat is found in 2:1 clay minerals The ideal full-cell chemical formula forkaolinite is Si4IVAl4VIO10(OH)8 and there is little, if any, isomorphicsubstitution in kaolinite

Other dioctahedral kaolins are dickite, nacrite, and halloysite (Table 2.4).The ideal full-cell chemical formula for these clays is the same as that for

FIGURE 2.5. Structural scheme of soil minerals based on octahedral and tetrahedral sheets From Schulze (1989), with permission.

kaolinite, with the major difference being the stacking sequence of layers.Halloysite (Fig 2.5) has water molecules between each 1:1 layer and theability to adsorb large quantities of monovalent cations such as NH4+ Dryingwill cause the water molecules to be removed and the clay layers to collapsetogether (Newman and Brown, 1987) Halloysite is also characterized by atubular morphology, whereas kaolinite, when examined with microscopictechniques, has a platy structure.

Two trioctahedral serpentines that are 1:1 clays are antigorite andchrysotile (Table 2.4) These clays have all three octahedral positions filledwith Mg There is also no substitution in either the tetrahedral or octahedralsheets, and thus they have no permanent negative charge

for the dioctahedral 2:1 type clays is Si8IVAl4VIO20(OH)4 (pyrophyllite) The two representative clays of the pyrophyllite–talc group are pyrophyllite,

a dioctahedral clay, and talc, a trioctahedral clay (Table 2.4) The layer charge

per half-cell formula unit (full-cell formula unit divided by 2), x, on these

clays is 0 due to no apparent isomorphous substitution In addition, there is

no interlayer space in pyrophyllite and talc (Newman and Brown, 1987).However, even though there is little if any permanent negative charge onpyrophyllite, the edge sites that are present can have a significant influence

on the retention of metals and on various physical properties of the clay(Keren and Sparks, 1994)

smectite–saponite group are characterized by a layer charge of 0.2–0.6 perhalf-cell formula unit (Table 2.4) The group includes the subgroupsdioctahedral smectites and trioctahedral saponites The dioctahedralsmectites are represented by montmorillonite, beidellite, and nontronite

The ideal half-cell chemical formula for montmorillonite is M0.33,

H2OAl1.67(Fe2+, Mg2+)0.33Si4O10(OH)2, where M refers to a metal cation in

the interlayer space between sheets The tetrahedral cations are Si4+and theoctahedral cations are Al3+, Fe2+, and Mg2+ One can calculate the negativecharge in the tetrahedral sheet as 0 and in the octahedral sheet as –0.33.Thus, the net negative charge is –0.33, which is balanced by exchangeable

cations represented by M0.33 The other feature that characterizesmontmorillonite is the presence of water molecules in the interlayer space(Fig 2.5) This causes montmorillonite to take on shrink–swell characteristics.The major difference between montmorillonite and the other twodioctahedral smectites, beidellite and nontronite, is that isomorphoussubstitution in these minerals occurs in the tetrahedral sheet (i.e., Al3+substitutes for Si4+) rather than in the octahedral sheet Nontronite also is anFe-bearing mineral with Fe3+ being the predominant element in theoctahedral sheet

The trioctahedral members of the smectite–saponite group are saponiteand hectorite In saponite, substitution occurs in the tetrahedral sheet In

hectorite, a Li-bearing mineral, substitution occurs in the octahedral sheet(Table 2.4).

vermiculites have a layer charge of 0.6–0.9 per half-cell formula unit.Dioctahedral vermiculite is characterized by substitution in both thetetrahedral and octahedral sheets, while trioctahedral vermiculite hassubstitution only in the tetrahedral sheet and all three of the octahedralcation positions are filled with Mg (Table 2.4)

While vermiculites are characterized by a platy morphology, similar tothat for micas, they contain interlayer water (Fig 2.5) Vermiculites formmainly from weathering of micas, particularly phlogopites and biotites, andinterlayer K+ is replaced by other interlayer cations However, vermiculiteshave a layer charge lower than that of micas, and the Fe2+in the octahedralsheet is oxidized to Fe3+(Newman and Brown, 1987)

Let us take the half-cell chemical formula for dioctahedral vermiculitegiven in Table 2.4, since there is substitution in both the tetrahedral and octahedral sheets, and see how one would calculate the net layer charge

on the clay:

M0.74H2O (Si3.564+ Al0.443+ ) (Al1.43+Mg0.32+Fe0.33+) O10(OH)2The total positive charge on the clay is Si4+(3.56 × 4+= 14.24) + Al3+(1.4 + 0.44 = 1.84 × 3+= 5.52) + Mg2+(0.30× 2+= 0.60) + Fe3+(0.30× 3+

This negative charge on the clay is balanced out by metal cations in the

interlayer space represented as M0.74in Table 2.4 Please note that since H2O

is in the interlayer space and is not structural water, it is not considered incalculating the net layer charge

2:1 Clays (Illite Group). Illite has a layer charge of about 0.8 per half-cell

formula unit, intermediate between smectite and mica Grim et al (1937)

developed the term illite to describe clay-size mica that was found inargillaceous rocks Other terms that have been used in lieu of illite arehydromica, hydromuscovite, hydrous illite, hydrous mica, K-mica,micaceous clay, and sericite Illite has more Si4+, Mg2+, and H2O but lesstetrahedral Al3+ and K+ and water than muscovite While K+ is thepredominant interlayer ion along with divalent ions such as Ca2+and Mg2+,

NH4+can also occur in illite

formula unit and are both dioctahedral, e.g., muscovite and paragonite, andtrioctahedral, e.g., biotite, phlogopite, and lepidolite (Table 2.4) With the

TABLE 2.4. Partial Classification of Phyllosilicate Clay Minerals

Composition (half-cell chemical formula unit)

Cations

(trioctahedral) Smectite–saponite Smectites Montmorillonite (M0.33, H2O)b Si4 Al1.67c(Fe 2+ ,Mg)0.33 O10(OH)2

Vermiculite Dioctahedral Dioctahedral (M0.74, H2O) (Si3.56Al0.44) (Al1.4Mg0.3Fe0.33+ ) O10(OH)2

vermiculite vermiculite Illite Illite Illite M0.74(predominately K) (Si3.4Al0.6) (Al1.53Fe0.22Fe0.03Mg0.28) O10(OH)2

(x > 0.6, x < 0.9)

brittle mica

y Soil Minerals

Composition (half-cell chemical formula unit)

Cations

exception of paragonite, a Na-bearing mica, the other micas have K+in theinterlayer positions to satisfy the negative charge resulting from isomorphoussubstitution Thus, micas are major K-bearing minerals in soils and as theyweather, the nonexchangeable K is released for plant uptake Weathering convertsmicas to partially expansible 2:1 clay minerals such as illite and vermiculite Thereleased K+ from layer and edge weathering results in frayed edges and

“wedge zones” (see Chapter 9 for a discussion of these) that play profoundroles in K selectivity and K fixation (Sparks, 1987)

since they are 2:1 clays with a hydroxide interlayer, either gibbsite-like[Al(OH)x] or brucite-like [Mg(OH)x ] where x is <3, that is continuous across

the interlayer sheet and is octahedrally coordinated This sheet is positivelycharged because there are fewer than 3 OH–per Al3+in the sheet The interlayersheet is bound to the 2:1 clay electrostatically, and the tetrahedral layer is bonded

to the interlayer sheet by hydrogen bonding

Chlorites can be trioctahedral in both octahedral sheets, i.e., theoctahedral sheet of the 2:1 layer and the interlayer octahedral hydroxidesheet, and are referred to as tri,trioctahedral chlorites Chlorites that aredioctahedral in the 2:1 layer and trioctahedral in the interlayer hydroxidesheet are referred to as di,trioctahedral chlorites

intermediates between smectites and vermiculites The interlayer space iscomposed of exchangeable cations and gibbsite-like or brucite-like islandsthat are not continuous as in chlorite Accordingly, the interlayer space doesnot collapse on heating as readily as with smectites and vermiculites but itdoes collapse more easily than the complete interlayer hydroxy sheets inchlorite A prominent intergrade clay mineral in many southeastern andmid-Atlantic United States soils is hydroxy-interlayered vermiculite (HIV),which is characterized by hydroxy–Al interlayers The hydroxy–Al in theseacid soils is an important source of nonexchangeable Al In alkaline soils theinterlayer material is hydroxy–Mg

INTERSTRATIFIED CLAY MINERALS

Since the 2:1 and 1:1 layers of clays are strongly bonded internally but areweakly bonded to each other, layers can stack together to form interstratifiedclays Examples include interstratified smectite with talc-type units andsmectite and mica units such as 1:1 mica–dioctahedral smectite or hectorite.Another example would be interstratification of smectite or vermiculite withchlorite For example, a 1:1 regularly interstratified chlorite–smectitestructure would contain four tetrahedral sheets, two in each of the 2:1 layers,three octahedral sheets, one in each of the 2:1 layers and the other in thehydroxidic interlayer, and one expanding interlayer space containing theexchangeable cations The anion content is O40(OH)20 Interstratification ofsmectite with kaolinite can also occur (Newman and Brown, 1987)